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DIAGNOSTIC STUDIES OF THE THERMAL STRUCTURE OF THE TROPICAL ATLANTIC OCEAN

J. Merle ORSTOM, Université de Paris, Paris, France

1. INTRODUCTION

1.1 T r o p i c a l A t l a n t i c ocean and c l i m a t e s t u d i e s

The t r o p i c a l r e g i o n s c o v e r a large p a r t o f t h e g lobe and p l a y a key r o l e i n t h e l a r g e scale i n t e r a c t i o n s between t h e atmosphere and t h e oceans , on a g l o b a l b a s i s . Changes i n SST c o n d i t i o n s i n t h e t r o p i c a l areas have been shown t o have a n i n f l u e n c e on t h e g l o b a l a tmospher ic c i r c u l a t i o n (HOREL and WALLACE - 1981 ; PAN and OORT- 1983) . A number of Genera l C i r c u l a t i o n Models (GCM) experiments led . t o t h e same c o n c l u s i o n s . The t r o p i c a l oceans p l a y a l s o a major role i n mer id iona l h e a t t r a n s p o r t (OORT and VONDER HAAR - 1976) . Furthermore t h e v i c i n i t y of t h e e q u a t o r a n d t h e van i sh ing of t h e C o r i o l i s paramameter imply t h a t a s t r a t i f i e d ocean can respond s t r o n g l y and r a p i d l y t o bas inwide wind f l u c t u a t i o n s and t h e r e f o r e . t h e wind f o r c i n g i s r e s p o n s i b l e f o r a l a r g e p a r t of t h e changes of t h e the rma l s t r u c t u r e and of t h e Sea S u r f a c e Temperature (SST) on a s e a s o n a l and i n t e r a n n u a l t i m e scale.

S e v e r a l s e q u e n t i a l s t u d i e s of t h e s t e a d y s t a t e and of t h e low-frequency v a r i a b i l i t y of t h e t r o p i c a l A t l a n t i c ocean have been c a r r i e d o u t du r in? t h e l as t decades (EQUALANT i n 1962-63, GATE i n 1974, FGGE i n 1979 and more r e c e n t l y SECTION and SEQUAL/FOCAL). The two l a s t exper iments had t h e o b j e c t i v e s of unders tanding t h e s e a s o n a l response of t h e g l o b a l t r o p i c a l A t l a n t i c b a s i n t o t h e s e a s o n a l l y va ry ing winds, T h i s s c i e n t i f i c i n t e r e s t f o r t h e A t l a n t i c ocean h a s many j u s t i f i c a t i o n s . Its direct c o n t r i b u t i o n t o t h e f o r c i n g of t h e g l o b a l a tmosphere is c e r t a i n l y weaker than t h a t of t h e Pacif ic and I n d i a n ocean , b u t it canno t be omi t t ed and fu r the rmore it r e p r e s e n t s and idea l case s t u d y of t r o p i c a l oceans dynamic ( l a b o r a t o r y ) . The t h r e e oceans are dynamica l ly s i m i l a r and of t h e t h r e e t h e A t l a n t i c ocean is t h e most p r e d i c t a b l e , p o t e n t i a l l y b e . c a u s e of its l i m i t e d z o n a l e x t e n s i o n and t h e r e l a t i v e s i m p l i c i t y of t h e wind f ie ld . A s a consequence, t h e r e sponse of t h e t r o p î c a l A t l a n t i c is almost i n phase wi th t h e wind changes, which are u s u a l l y v e r y r e g u l a r and dominated by t h e annual s i g n a l , excep t p a r t i c u l a r y e a r s , when a b r u p t wind b u r s t s e x c i t e e q u a t o r i a l waves t h a t are b e l i e v e d t o have a r e l a t i o n w i t h t h e observed warm and c o l d SST anomal i e s i n t h e Gulf of Guinea (SERVAIN, PICAUT and MERLE - 1982 i SERVAIN, PICAUT and BUSALACCHI - 1985) .

Th i s l o g i s t i c a l l y compact t ropical A t l a n t i c ocean h a s a number o f o t h e r v i r t u e s . It is m o r e accessible and i n s e v e r a l r e s p e c t s b e t t e r known than t h e o t h e r t r o p i c a l oceans . The p a s t experiments mentionned above, and p a r t i c u l a r l y SEQUAL/FOCAL, have provided what can be cons ide red as t h e best t r o p i c a l basinwide d a t a set documenting t h e l o w f r equency r e s p o n s e of a g l o b a l ocean t o t h e wind f o r c i n g . A number of A f r i c a n and South-American c o u n t r i e s are s u f f e r i n g s e v e r e climatic c o n d i t i o n s ( C h i l i a n , North-Brazi l and Sahe l d r o u g h t s ) which can be p a r t i a l l y c o n s i d e r e d as t h e r e s u l t of p a r t i c u l a r t h e r n a l c o n d i t i o n s of t h e t r o D i c a l A t l a n t i c ocean anä s o s s i b l y cou ld be f o r e c a s t e d a t s h o r t t i m e scales (NOURRA anä

Fonds Documentaire ORSTOM

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S'n'UicLX - 1981). Finally t h e Atlantic o c e a n i.: a meridional ocean ar12 i t s t r o p i c a i part is one of the ?;rLnci?al meridional heat exchanae reaicn of the world's oceans and so has ail important influence on t h e global earth's climacz at lonquer time scales (F'iq.11.

FIG.l The ocean rean heat transFort irì 10 13 W fro= SC:??EZ/iC;79>

1.2 Two forcings and two responses of the tropical oceans.

A thin thermocline layer is almost permanent in the tropics. It

cold waters in an almost perfect two-layers ocean. This thermal structure leads to.separate the response of the tropical ocean into two parameters : (i) the depth of the thermocline, (ii) the temperature of the mixed layer (or of the Sea Surface). These parameters are, at the first order, independents. The depth of the thermocline i s the dynamical response of the ocean to the wind forcing (either local, remote or global). The temperature of the mixed layer or SST is the result of the local thermodynamic forcing of the atmosphere. These simplified views apply especially in the western regions where the thermocline is deep. In the Eastern equatorial regions where the thermocline is shallow and reaches Che surface it creates particular surface conditions (cold water) that in turn modify the conditions of the local thermodynamic forcing (the ocean absorbs heat from the atmosphere). And there the two forcings and the two corresponding responses of tihe ocean are clearly not independent.

Nevertheless understanding the response of the ocean in the tropics require, in a first approach, to study the oceanic dynamical. response (depth of the thermocline) to the wind forcing. The second order phenomenon is the thermodynamic coupling (forcing and response) that is almost entirely confined in the oceanic mixed layer (.

We propose a synthetic overview of the present knowledge on the large scale, low frequencies forcings and responses of the tropical Atlantic ocean. In chapter two the forcings over the tropical Atlantic are described. In chapter three, the dynamical response of the ocean. In chapter four, the heat budget and heat transport. In chapter five, the SST response. In chapter six, conclusions are derived.

- separates a well mixed and warm superficial layer from the deeper

..

. v - 25

2. THE FORCINC-S

Seve ra l n i l l i o n s u r f a c e o b s e r v a t i o n s from merchant a n d weather s h i p s , compiled by t h e N a t i o n a l C l i m a t i c C e n t e r and ëes iganced TDF-II, have been used by many a u t h o r s t o comgute varkous me teo ro log ica l e lements o f t h e s u r f a c e oceans .

I n t h e t r o p i c a l A t l a n t i c HASTEMRATH and LAKB ( 1 9 7 7 ) produced a c l i m a t o l o g i c a l A t l a s i n c l u d i n q monthly c h a r t s of r e s u l t a n t wind, d ivergence , wind s t e a d i n e s s , w i n d - s t r e s s c u r l , r e l a t i v e v o r t i c i t y . HELLERMAN ( 1 9 8 0 ) computed t h e annua l mean and s t a n d a r d d e v i a t i o n s of wind-s t ress , w ind- s t r e s s c u r l s and d ive rgence of Ekman t r a n s p o r t i n t h e t r o p i c a l A t l a n t i c . BUNKER ( 1 9 7 6 ) and HASTENRATH and LAMB (1978) a l s o c a l c u l a t e d t h e i n d i v i d u a l e n e r g y f l u x e s a c r o s s t h e sur face t o o b t a i n monthly and annua l maps of t h e s e parameters . They a l s o comDuted t h e r e s u l t i n g n e t ene rgy f l u x across t h e s u r f a c e .

2.1 Wind f o r c i n g

The wind stress means fo r March and August are p r e s e n t e d Fig.2 from HELLERMAN (1980). The n o r t h e a s t and s o u t h e a s t t radewinds ar2 Seen to p r e v a i l o v e r t h e wes te rn p a r t o f t h e A t l a n t i c ocean. Alona t h e AfricaR c o a s t and i n t h e Gulf of Guinerl, t h e x inds are predominant ly mer id iona l . The t radewinds converqe on t h e

FIG. 2 Mean Wind-stress for the months of March and August (from HELLERMAN - 1979)

I n t e r t r o p i c a l Convergence Zone (ITCZ) which p r e s e n t a l a r g e s e a s o n a l migrat ion. I n March it i s p r a c t i c a l l y a l o n g t h e equa to r , i n August it i s i n i ts nor the rnmos t p o s i t i o n from 1 0 0 ~ i n t h e w e s t t o 15ON i n t h e E a s t .

The seasona l and i n t e r a n n u a l v a r i a b i l i t y of t h e zona l and m e r i d i o n n a l component of t h e wind- s t r e s s have been s t u d i e d more r e c e n t l y by SERVAIN, PICAUT and BUSALACCHI (1985). They u s e a s i x t e e n years SST and wind data set monthly ave raged from January 1964 t o december 1979 i s s u e d f r o m t h e TDF-II data f i l e . Regions of t h e m a x i m u m s e a s o n a l v a r i a b i l i t y of t h e z o n a l and mer id iona l wind stress (Fig.3 upper p a n e l s ) are c o n t a i n e d w i t h i n an envelope d e f i n e d by t h e s e a s o n a l e x c u r s i o n of t h e ITCZ. The m a x i m a of t h e zona l wind stress f l u c t u a t i o n s straddle t h e mean p o s i t i o n of t h e ITCZ whereas t h e largest d e v i a t i o n s o f the m e r i d i o n a l wind stress occur along t h i s l i n e . I n t e r a n n u a l wind stress f l u c t u a t i o n s do n o t n e c e s s a r i l y c o i n c i d e w i t h t h e regions o f t h e maximum s e a s o n a l v a r i a b i l i t y . The largest y e a r t o y e a r changes i n t h e zona l wind stress are a t t h e poleward ex t r emes of t h e t r o p i c a l area. The ampl i tude of t h e i n t e r a n n u a l v a r i a b i l i t y of t h e mer id iona l wind stress i s n e a r l y c o n s t a n t . Minima are however n o t e d i n s i d e t h e gul f of Guinea and o f f s h o r e t h e coasts o f Guiana and Br 'az i l (F ig .3 lower pane l s ) .

\

V - 26

INTERANNUAL VARIABILITY OF fi

FIG. 3 Seasonal and interannual-variability of t h e zonal and meridional component of wind-stress(from SERVAIN e t a l - 1985)

2.2 H e a t f l u x f o r c i n g .

The energy exchange a t t h e ocean s u r f a c e h a s been computed f irst by Budyko (1963) from t h e equa t ion :

A = Q ( 1 -o() - I - L - S

where A i s t h e monthly h e a t g a i n by t h e ocean , Q (1 -4) is t h e solar r a d i a t i o n absorbed by t h e o c e a n d i s t h e sur face a lbedo , I i s t h e n e t i n f r a r e d r a d i a t i o n exhange a t t h e s u r f a c e , L i s t h e l a t e n t h e a t f l u x due t o eyapora t ion and S is t h e s e n s i b l e h e a t exchange w i t h t h e atmosphere. BUNKER (1976) and HASTENRATH and LAMB (1978) u s e d t h e same g e n e r a l formula wi th some d i f f e r e n c e i n t h e computa t ion of each component of t h e budget, T h e i r r e s u l t s are n o t d i f f e r e n t i n t h e i r g e n e r a l p a t t e r n S . b u t d i f f e r s i n t h e i r ex t remes v a l u e of about 10 %. For t h e t r o p i c t h e m o s t d e t a i l e d p i c t u r e are t h o s e of HASTENRATH and LAMB (1978) and HASTENRATH (1977) . Fig.4 p r e s e n t s t h e annual mean of t h e n e t h e a t f l u x a t t h e s u r f a c e . The cold w a t e r s o f t h e e q u a t o r i a l and coastal upwel l ing regions r educe t h e l a t e n t h e a t loss by weaker evapora t ion and t h e r e f o r e t r ap t h e h e a t f r o m t h e atmosphere. Seasona l v a r i a t i o n s of t h e n e t s u r f a c e h e a t f l u x broadly f o l l o w t h e annual c y c l e of i n s o l a t i o n w i t h l a r g e s t g a i n i n t h e r e s p e c t i v e summer and loss i n t h e w i n t e r h a l f - y e a r (F ig .5 ) . The maximum ocean h e a t gai.n appears n e a r t h e

t h e m o s t i n t e n s e upwell ing pe r iod . The h i g h e s t v a l u e s (120 W/m 1 are found i n August nea r I O o W (HASTENRATH and LAMB'1978).

e q u a t o r (between O and S O S ) i n no r the rn summer and co r re spon 9 to

v - 27

FIG. 4 Mean a n n u a l n e t h e a t f l u x a t t h e air-sea i n t e r f a c e . Shaded area i n d i c a t e s

h e a t g a i n by t h e ocean .uni t s W/m2. (From HASTENRATH - 1977)

Y 1 1 1 1 , I I I I # ,

J F M A M J ' J A S O N D '

FIG. 5 Seasonal v a r i a b i l i t y of t h e n e t h e a t f l u x zonaly averaged i n t h e t r o p i c a l

A t l a n t i c from 300N t o 30"s. Shaded area - . i n d i c a t e s h e a t g a i n by t h e ocean. U n i t s W/mZ (From HASTENRATH - 1977)

3. THE DYNAMICAL RESPONSE

3.1 Observations

A merged data set including MBT, XBT, and NANSEN - CTD observations, covering the 2OoS - 30°N tropical Atlantic ocean until 1978, has been used to study the seasonal changes in the upper ocean thermal structure (MERLE 1983).

Fig.6 shows the topography of the 2OoC isothern in :+larch and August. The difference is clear. In March the 2OoC isotherm present a simple structure with a zonal East-west general slcpe and a slight upwelling along the equator. In August the pattern is different. A zonal downwelling appear at about 3-4ON boardering further North a zonal ridge under the northernmost position of the ITCZ (lOON - lSON). The maximum seasonal changes are-found in three regions : (i)lOo North (ii) equator and West (iii) Eastern equator (Fig.7 upper left panel). Annual signals of thermocline depth variations are not in phase in the three regions (Fig.7 upper right panel). The maximum depth of the thermocline is observed in fall in the West equatorial region and in Spring in the Eastern equatorial region. A double seasonal tilt of the thermocline is resulting of these phase changes : one along the equatorial plan, the other one along a zonal pivot line situated under the mean position of the ITCZ. This double axis of rotation of the thermocline on a seasonal time scale is the result of two distinct responses of the ocean to the wind forcing : (i) a remote equatorial response which mainly affects the Eastern equatorial region. (ii) A local North equatorial response associated to the seasogal migration of the ITCZ that induces a change in the sign and the magnitude of the curl of the wind stress.

The Equatorial response has been first tentati'vely explained by MOORE et al (1978) that proposed a simple remote forcing mecanism to account for the Guinea Gulf upwelling. According to this hypothesis an increase of the easterly wind in the western

V - 28

-10"

s aoo . 60"

FIG. 6 Topography of the 20°C isotherm in March and August.

I -

I -

equatorial Atlantic excites an internal equatorial KELVIN wave that propagate Eastward and affect the thermal skructure in the Guinea Gulf. Latter on,.observations of HOUGHTON (1983) in the Guinea Gulf confirmed that a trapped equatorial phenomenon affect the thermocline near ~ 4 O W (Fig.8). Furthermore SERVAIN et al (1982) found a hight correlation -beetween the wind stress anomalies off Brazilian coast and the SST anomalies in the Guinea Gulf one month later (see chapter 5 ) .

The north equatorial response has been studied by GARZOLI and KATZ (1983). They found that the seasonal reversal of the North Equatorial countercurrent and the associated change of the depth of thermocline beetween 3ON and 10°N (Fig.9) is due to the changes of the wind-stress curl throught the combined mechanisms of the local EKMAN pumping and the divergence of the geostrophic currents.

3.2 Modelling

Important progresses have been achieved during the last decade on modelling the tropical oceans. Historically, theoretical approaches have evolved in several stages. First idealized winds blowing over shallow waters models were used both in the Pacific and in the Atlantic (Mc CREARY 1976 ; CANE and SAEiACHICK - 1981 ; BUSALACCHI and O'BRIEN - 1980 ; O'BRIEN, ADAMEC and MOORE - 1978) to understand the forced ocean response to abrupt changes of the wind forcing. The equatorial Kelvin waves and their effect on the thermal structure of the eastern equatorial regions have been predictea by these theories before being observed in the Pacific and the Atlantic (KNOX and HALPERN - 1982 ; SEQUAL group - personal communication), Then the same shallow water models were used with realistic climatologic winds to simulate the mean seasonal response of the ocean (BUSALACCHI and O'BRIEN - 1981 ; BUS'ALACCHI and PICAUT - 1983 ; DU PENHOAT and TREGUIER - 1984). Such modelling results for the Atlantic are presented Fig.7, Both the amplitude and phase of the thermocline depth seasonal variability are correctly

40 30 20 10 w O 40 E I O SO 40 30 20 i o w O 40 F ._

Fig. 7 Upper panels : amplitudes (left) and phase (right) of the seasonal varia- tion of the depth of the thermocline (20T isotherm) Intermediate panels : Amplitude and phase of the seasonal variation of the depth of the surface separation in a 2 layers model forced by climato- logical winds (From BUSALACCHI and PICAUT - 1983). Lower panels : Amplitude in a two layers model. forced by climatological winds (From DU PENHOAT and TREGUIER - 1984).

CL-

and phase of the surface dynamical topography

e I

N .o

V - 30 simulated. more recently, realistic winds and surface thermodynamic f l u x e s have been used to force sophisticated three-dimensional models (PHILANDER and PACANOWSKI - 1084) that simulate most of the known features of the seasonal variabilit-i of the troyical Atlantic ocean and parricularly the surface currents (Fig.10).

2**-

- . . .

Isotherm Displacement im) at 4"W Apr - Aug

- 1 1 1 1 1 1 I ' , 1 I I 1

FIG. 8 Isotherm displacement beetween April. and August a t 4OW i n the Guinea Gu l f (From HOUGHTON - 1983)

50 40' 30 20

L ONG/TUD€ FIG. 9 Seasonal longitudes changes of the depth of the thermocline. from 3ON to

7ON (From GARZOLI and KATZ - 1983) compared with the surface currents chan- ges obtained by ship-drift (from RICHARDSON - 1984)

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FIG. 10 Surface currents in the tropical Atlantic Ocean, February and August, as simulated with a multi-level numerical model (From PHILANDER and PACANOWSKI 1984)

4 . HEAT BUDGET AND HEAT TRANSPORT

Seasonal variations of heat content follow the seasonal variations of the thermocline depth (Fig.11 a and b). The 0-300 meters. heat content shows a large maximum of amplitude in the Wgst and2 at the North of the equator with values close to 6.10" J/m . An gaste n equatorial maximum is also observed with values over 10.10 J/m at the African coast close to the equator. The phase map shows also a phase difference between the Eastern and the Western part of the equatorial band with a transition point at about 1 S O W . Another change phase is observed along a zonal line near 6-80N which is similar to the one described for the 2OoC isotherm depth and situated under the mean position of the ITCZ.

These differences in the seasonal variation of heat content are also marked in the heat storage ( S o ) that is the difference of the time rate of heat content change (AHC) minus the net heat flux exchanged at the air - sea interface (FI

5

SO 4 H C - F

V - 3 2

FIG. 11 a and b,amplitude and phase of the seasonal variation of heat content in the layers 0-300 meters. '

I n t h e near E q u a t o r i a l band (6ON - 6OS), t h e s e a s o n a l v a r i a t i o n , of t h e t i m e rate of h e a t c o n t e n t is about 10 times l a r g e r t h a n ' t h e . s e a s o n a l v a r i a t i o n of t h e n e t h e a t f l u x through t h e s u r f a c e g iven by HASTENRATH and LAMB (1978) (Figs.12a-12b). I n t h e wes te rn r e g i o n . a d ivergence (or e x p o r t ) of h e a t i s observed f r o m October t o May wi th a maximum i n March ; a h e a t convergence (o r fmpor t ) o c c u r s f r o m June t o September i n t h e 6ON-6OS r e g i o n as a whole. I n t h e eastern r eg ion (Gulf of Guinea] h e a t d ivergence occur s i n t w o p e r i o d s : t h e main p e r i o d is April-August and t h e secondary p e r i o d i s November-December, Thus i n an annual mean t h e e q u a t o r i a l A t l a n t i c r eg ion is e x p o r t i n g h e a t b u t n o t th roughout t h e y e a r (MERLE 1980 a ) .

Furthermore t h e s e zona l and mer id iona l phase changes in h e a t storage i m p l y t h a t h e a t is s l o s h e d Eas t -West b u t also r e d i s t r i b u t e d mer iona l ly . Fig.l.3a and b p r e s e n t t h e t i m e rate change of h e a t c o n t e n t for t w o months ( June and September) t h a t f r a m e t h e summer e q u a t o r i a l upwell ing i n t h e Gulf of Guinea. h e a t convergence (d ive rgence ) appears a t a rate of t h e o r d e r of 300 watts p e r m e t e r square.

LAMB and BUNKER (1982) computing t h e 0-500 m e t e r s h e a t c o n t e n t and u s i n g t h e n e t h e a t s u r f a c e f l u x provided by BUNKER (1976) i n t e g r a t e d t h e mer iona l component of t h e h e a t t r a n s p o r t th roughout t h e A t l a n t i c from 70°N t o 20°S. H e fpynd an annual mean mer id iona l h e a t t r a n s p o r t of t h e o r d e r of 150.10 compat ib le with prev ious e s t i m a t i o n of EMIG (19671, STOMMEL (1979) and a l so BRYDEN and HALL. (1980) t h a t used d i r e c t hydrographic o b s e r v a t i o n s along t h e 27"N p a r a l l e l .

W i n t h e t r o p i c a l A t l a n t i c

v - 3 3

A b

F IG. 12a Annual variation o f time rate of change of heat content (interrupted l ine ) and annual variation of net heat oceanic gain ( f u l l l i ne ) i n the western equatorial Atlantic region. Units are W/mZ. Dashed areas represent diver-

A s i n Fig.10 a, except f o r the eastern region (From MERLE - 1980 a)

gence o f heat.

F IG. 12b

F I G . 13a Rate of change of heat content from O to 300 Meters i n June ( i n W/m')

F IG. 13b A s in Fig.1 a f o r September.

v - 3 4

Furthermore LAMB and BUNKER p r e s e n t e d a s e a s o n a l var ic l t ion of t he mer id ionnal h e a t t r a n s p o r t i n d i c a t i n g a s u r p r i s i n g r e v e r s e of t h e h e a t t r a n s p o r t f lowing southward i n f a l l a l l a long t h e A t l a n t i c f rom 70"N t o 20"s (Fiq.14).

7 u N 7U N

66 eu

SQ

UT

I)' XT

to

lb

U U

10

X I ZVS J F Y A Y J , J A 5 O N O 1

F- So = D I V(T 1 TY F

FIG. 14 Seasonal variation of zonal integration i n the Atlantic of : - Net heat flux a t the air-sea interface (F) - Divergence of heat transport (T), So being the time ra te change of heat content i n layer 0-500 meters. - Meridionnal component of.heat transport i n the layer 0-500 meters (Ty). (From LAME - 1983).

5. THE SZA SURFACE TEMPERATURE RESPONSE

5.1 Observations of sea S u r f a c e t empera tu re v a r i a b i l i t y .

The s l o p i n g of t h e t h e r m o c l i n e which rises f r o m w e s t t o east i s p a r t l y r e s p o n s i b l e for t h e c o o l e s t waters t h a t are found i n t h e e a s t e r n t r o D i c a l A t l a n t i c , The s p a t i a l s t r u c t u r e of t h e ampl i tude of s e a s o n a l SST v a r i a b i l i t y ressembles t h e mean s t a t e (Fig.15 - SERVAIN e t a l . - 1985). Regions wi th t h e Greatest s e a s o n a l v a r i a b i l i t y c o i n c i d e w i t h r e g i o n s of low SST. The sha l low thermocl ine i n t h e e-astern trosical A t l a n t i c t o g e t h e r w i t h v e r t i c a l mixing induces maximum s e a s o n a l SST v a r i a b i l i t y ( U - = 1.4-3.4OC) i n t h e s e a s o n a l upwel l ing zones a l o n g t h e coasts of Maur i t an ia and Seneqal (NE7 A f r i c a ) , t h e n o r t h e r n and s o u t h e r n coasts of t h e Gulf of gu inea , and a long t h e e q u a t o r n e a r 1OOW. The l o c a t i o n of t h e minimum SST v a r i a b i l i t y (os0 .4OC) is c o i n c i d e n t w i t h t h e thermal e q u a t o r and t h e mean I n t e r t r o p i c a l Convergence Zone ( ITCZ) .

maximum (o- = 0.6 -1.OOC) i n t h e c o a s t a l and equatorial upwel l ing r e g i o n s b u t minimum o f f s h o r e t h e n o r t h e r n c o a s t of B r a z i l .

The i n t e r a n n u a l v a r i a b i l i t y of SST, as d e p i c t e d i n Fig.15 i s

5.2 A r e m o t e f o r c i n g mecanism t o e x p l a i n t h e SST v a r i a b i l i t y i n t h e Gulf o f Guinea ?

MOORE e t a l (1978) s u g g e s t e d t h a t remote f o r c i n g by winds i n t h e wes tern At lan t ic may p r o v i d e an e x p l a n a t i o n of t h e SST v a r i a b i l i t y i n t h e G u l f of Guinea. - .

v - 35 INTERANNUAL VARIPBILITY CIF SST SEASONAL VARIABILITY OF SST

20H

O

10s

mn

O

10s

FIG. 15 Seasonal and Interannual variabil i ty of SST (From SERVAIN et al - 19E?- An a n a l y s i s of SST and s u r f a c e winds i n selected areas of t h e

t r o p i c a l A t l a n t i c by SERVAIN e t a l (1982) i n d i c a t e d t h a t t h e nonseasonal v a r i a b i l i t y of SST i n t h e e a s t e r n e q u a t o r i a l A t l a n t i c (Gulf o f G u i n e a ) i s h i g h l y c o r r e l a t e d wi th t h e noiiseasonal v a r i a b i l i t y of t h e zonal wind stress i n t h e w e s t e r n e q u a t o r i a l A t l a n t i c (F ig .16) . A nega t ive ( p o s i t i v e ) anomaly of t h e wind stress n e a r t h e n o r t h B r a z i l i a n c o a s t i s f o l l o w by a p o s i t i v e ( n e g a t i v e ) anomaly i n t h e Gulf of Guinea about one month l a t e r . Furthermore, t h e correlation between t h e l o c a l winä stress ahomaly and SST anomaly of SST i n t h e Gulf of Guinea is c o n s i d e r a b l y s m a l l e r . These r e s u l t s i n d i c a t e t h a t remote f o r c i n g i n t h e wes te rn e q u a t o r i a l A t l a n t i c ocean i s a f f e c t i n g t h e e a s t e r n e q u a t o r i a l A t l a n t i c sea s u r f a c e tempera ture .

Some p a r t i c u l a r years l i k e 1967 o r 1968 impu l s ive wind stress f l u c t u a t i o n s i n t h e Western e q u a t o r i a l A t l a n t i c c o u l d e x c i t e e q u a t o r i a l l y t r apped waves t h a t would induce SST f o r c e d r e sponse i n t h e Gulf of Guinea. The warming i n t h e Gulf of Guinea o f m o r e t h a n 2OC i n July-August 1968 w a s preceded by abnormal wind stress f o r c i n g o f f t h e nor thern B r a z i l coast one month b e f o r e ( F i g . 171. These w a r m e v e n t s r e l a t e d t o abnormal wind s i m i l a r t o t h e "EL NINO" phenomenon i n t h e Eas te rn E q u a t o r i a l P a c i f i c have been called " A t l a n t i c EL NINO" (HISARD - 1980 ; MERLE - 1980b) . I n 1984 a w a r m e v e n t of unusual amplitude a f f e c t e d t h e E a s t e r n t r o p i c a l A t l a n t i c d u r i n g t h e r e l a x a t i o n phase of t h e 1982-83 P a c i f i c "EL NINO" . Some

, p o s s i b l e connexions beetween t h e two e v e n t s c o u l d be sea rched through t h e g l o b a l long t e r m p e r t u r b a t i o n of t h e t r a d e winds du r ing t h e p e r i o d 1982-1984.

V - 36

FIG. 17 .Comparison of monthly anomalies of SST i n GW area( dotted l i n e ) and TX i n BR area ( so l id l i n e ). From SERVAIN e t a l (1983)

6 - CONCLUSIONS

Three t i m e s s m a l l e r t h a n t h e t r o p i c a l P a c i f i c ocean t h e t r o p i c a l A t l a n t i c is a remarquable e q u a t o r i a l " l a b o r a t o r y " f o r s t u d y i n g t h e basic phys ic s of t h e coupled ocean-atmosphere system. A s m a l l ocean b a s i n h a s a s h o r t e r ad jus tmen t t i m e t h a n a l a r g e b a s i n , because t h i s ad jus tmen t t i m e i s p r o p o r t i o n a l t o t h e t i m e it t a k e s p l a n e t a r y waves t o p ropaga te a c r o s s t h e b a s i n . Thus, i n s t e a d of t h e Pac i f ic , t h e t r o p i c a l A t l a n t i c h a s almost no memory. The zona l s l o p e of t h e thermocl ine a l o n g t h e e q u a t o r , t h e d e n s i t y f i e l d and t h e i n t e n s i t y of t h e equator ia l c u r r e n t s v a r y a lmost i n phase w i t h s e a s o n a l changes i n t h e wind f i e l d . Even o u t s i d e of t h e e q u a t o r i a l band a n e a r e q u i l i b r i u m re sponse of t h e t r o p i c a l ocean i s observed. The s e a s o n a l reversal of t h e North E q u a t o r i a l c o u n t e r c u r r e n t i s i n phase w i t h t h e s e a s o n a l excur s ion of t h e ITCZ and t h e a s s o c i e t e d wind changes (G.UZOL1 anä KATZ - 1983). 3ecause of i t s s rca l le r s i z e t h e t r o p i c a l A t l a n t i c shows s e a s o n a l chanaes t h a t a r e dynamical ly s imi l a r t o i n t e r a n n u a i changes i n c.he F a c i f i c and make it easier t o observe.

The a v a i l a b l e d a t a and t h e models s i m u l a t i o n s have provided remarkable cohe ren t s p i c t u r e s o f t h e s e a s o n a l response of t h e t r o p i c a l A t l a n t i c ocean t o t h e wind f o r c i n g (F ig .7 ) . The double o s c i l l a t i o n of t h e the rmoc l ine on an E a s t - W e s t and North-South r o t a t i o n imply s e a s o n a l r e d i s t r i b u t i o n s o f h e a t z o n a l l y and mer iona l ly t h a t are an order of magnitude l a r g e r and o u t of phase w i t h t h e h e a t f l u x across t h e s u r f a c e (MERLE - 1980a) . T h i s means t h a t t h e dynamical i n t e r i o r r e sponse of t h e t rop ica l ocean is t h e m o s t impor tan t f o r unde r s t and ing and p r e d i c t i n g t h e ocean ' s r o l e i n t h e o v e r a l l climate system. Zonal e q u a t o r i a l h e a t t r a n s p o r t may account for "EL NINO" l i k e phenomenons on s e a s o n a l t i m e scale or i n t e r a n n u a l t i m e scale (MERLE - 1980b ; HISARD - 1 9 8 0 ) l i k e i n t h e P a c i f i c .

If a t t h e s e a s o n a l and i n t e r a n n u a l t i m e scales t h e A t l a n t i c ocean has a weaker i n f l u e n c e on t h e world climate than t h e P a c i f i c and I n d i a n oceans , it may have a dominant i n f l u e n c e on t i m e scales of decades o r c e n t u r i e s . The A t l a n t i c ocean is a .mer id iona1 ocean f u n c t i o n i n g as a h e a t p i p e , d e r i v i n g h e a t t o t h e North A t l a n t i c from t h e world ocean h e a t c i r c u l a t i o n sys t em (Fig.1). The narrow t r o p i c a l A t l a n t i c c o n c e n t r a t e a huge c r o s s e q u a t o r i a l merional h e a t f l u x and may of fe r t h e b e s t p l a c e i n t h e w o r l d t o observe and unders tand t h e oceanic h e a t t r a n s p o r t and i t s i n f l u e n c e on c l ima te .

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